Advent of Strong South Asian Monsoon by 20 Million Years Ago

Gregory J. Retallack,1,* Sunil Bajpai,2 Xiuming Liu,3

Vivesh Vir Kapur,2 and Santosh Kumar Pandey2

1. Department of Geological Sciences, University of Oregon, Eugene, Oregon 97403-1272, USA; 2. Birbal SahniInstitute of Palaeobotany, 53 University Road, Lucknow 226007, India; 3. School of Geographical

Sciences, Fujian Normal University, Fuzhou 350007, China

AB STRACT

The monsoonal paleoclimate of India has been critical for understanding the tectonic history of Himalayan and Ti-betan uplift over the past 60My.Monsoonal circulation in deep time has been inferred from variation in stable isotopesof tooth enamel, diatom blooms, and dust influx in the Indian Ocean and the advent of C4 grasses, but these proxies arecompromised by temperature, biotic, and source effects. Our study uses a proxy of carbonate distribution within pa-leosol profiles to infer appearance ofmonsoonal circulation ofmodern strength in theHimachal Pradesh segment of theHimalayan foreland by at least 20My ago, cued to HighHimalayan deformation and ongoing Tibetan Plateau uplift andretreat of the Paratethys Sea. Paleosol records also demonstrate declining chemical weathering with Himalayan andTibetan uplift, which was a force for global warming, rather than cooling, over the past 20 My.

Online enhancements: supplemental tables.

Introduction

Announced with thunder and lightning, the sum-mermonsoon of India is an eagerly awaited reprievefrom stifling heat, celebrated in Indian poetry backinto the Vedic Period of 3500 y ago (Murty 2014).Monsoon rains rouse parched soils to life and signalthe planting season for rice in India, Southeast Asia,and China. Monsoon winds from tropical seas bringsummer rain and are followed by monsoon windsfrom cold and dry northern lands in winter andspring. The intensity of the SouthAsianmonsoon isthe difference between dry-season and wet-seasonwind, temperature, and precipitation (Huber andGoldner 2012; Wang and Chen 2014). These sea-sonal differences are most extreme over the Indo-Gangetic plains of India, because of strong land-seatemperature gradients created by the Himalaya andthe Tibetan Plateau (An et al. 2001; Clift et al. 2008;Wang and Chen 2014), although opinion is dividedon whether the range (Boos and Kuang 2010) or the

plateau (Wuet al. 2012) ismore important.Theupliftof Asian mountains also has been considered im-portantforglobalweatheringandcooling (RaymoandRuddiman 1992), although the causation and direc-tionof thosechangesalsoaredisputed (Jacobsonetal.2002; Quade et al. 2003; Retallack 2013). Both con-troversies are addressed here fromnew evidence of pa-leoprecipitation from paleosols in India, Nepal, andChina.The link between uplift and monsoons is so

striking (An et al. 2001; Clift et al. 2008; Wang andChen 2014) that different ages of uplift have beenderived from different proxies for the advent of mon-soonal paleoclimate: late Eocene (40 Ma), from oxy-gen isotopic variation in growth rings of fossil snailsand mammals (Licht et al. 2014b) and from fossilwood growth rings (Licht et al. 2015); early Miocene(23 Ma), from geochemical and mineralogical indi-ces of chemical weathering (Clift et al. 2008); mid-Miocene (12.9 Ma), from particulate organic matterin the IndianOcean (Betzler et al. 2016); lateMiocene(10 Ma), from diatom blooms in the Indian Ocean(Gupta et al. 2004); Mio-Pliocene (7–5 Ma), from theadvent of C4 grasslands in South Asia (Quade et al.

Manuscript received February 13, 2017; accepted Septem-ber 5, 2017; electronically published November 29, 2017.

* Author for correspondence; e-mail: gregr@uoregon.edu.

1

[The Journal of Geology, 2018, volume 126, p. 1–24] q 2017 by The University of Chicago.All rights reserved. 0022-1376/2018/12601-0001$15.00. DOI: 10.1086/694766

1989); and early Pliocene (5 Ma), from dust influx indeep sea cores (An et al. 2001). These indirect proxiesmay be compromised by competing effects: oceanicversus continental sources of moisture and temper-ature versus precipitation for oxygen isotopic com-position (Licht et al. 2014b); nonseasonal versusseasonal climatic factors for chemical-weatheringindices of sediments (Clift et al. 2008) and dust dis-persal (An et al. 2001); and evolutionary innova-tion versus population explosion of phytoplankton(Gupta et al. 2004; Betzler et al. 2016) and grasslands(Quade et al. 1989, 1995). Our study brings to thequestion of monsoon antiquity new observationsof long sequences of paleosols in India, Nepal, andChina (fig. 1).

Pedogenic carbonate is a conspicuous feature ofmany paleosols (figs. 2, 3) widely used in paleoen-vironmental studies (Quade et al. 1989, 1995; Retal-lack 2005). Commonly, it is low-magnesium, mi-critic, replacive nodules that form below the soilsurface as soils dry out and respired soil CO2 de-clines at the end of the growing season (Breeckerand Retallack 2014). Depth to the horizon of carbon-ate nodules (calcic or Bk horizon in soil terminology)is thus related to both soil respiration in parts CO2per million (Breecker and Retallack 2014) and meanannual precipitation in millimeters, which is the cli-matic driver of biological productivity (Retallack 2005).Soils of desert regions have shallow calcic horizonsas a result of low productivity and precipitation, butsoils of subhumid regions have deep calcic horizonsreflecting high productivity and precipitation.

Indian monsoonal soils suffer both extremes: lowbiological productivity and precipitation for the hotdry season versus high productivity and precipita-tion for the warm wet season. Thickness of a soilwith carbonate nodules (measured as the distancebetween the highest and lowest nodules) is a usefulproxy for the difference in mean monthly precipita-tion between the driest and wettest months (Retal-lack 2005). For soils of the Indo-Gangetic plains ofIndia, there ismore than 100mmdifference inmeanmonthly precipitation. In addition, monsoonal soilsof India also have both shallow and deeply reachingroot traces and generally small nodules, with evi-dence of dissolution and reprecipitation of carbon-ate, and concentric laminae of goethite and hematite(Retallack 1991, 1995). Such nodules are not foundin summer-dry soils or paleosols, which have sili-ceous rhizoconcretions (Retallack 2004). Additionalevidence against summer-dry (Mediterranean) cli-matic seasonality for SouthAsia is pedogenic carbon-ate evidence for the advent during the late Mioceneof C4 grasses, which are still unknown in summer-dry climates (Quade et al. 1989, 1995). This study

thus focused on long time series of Miocene to Pleis-tocene paleosols in the Indo-Gangetic plains and Eo-cene to Miocene paleosols in the Lesser Himalaya.Oligocene to Pleistocene paleosol sequences were ex-amined directly north of the Tibetan Plateau in Chi-na’s Gansu Province, at the arid fringes of the EastAsianmonsoon system (Molnar et al. 2010; Sha et al.2015; Spicer et al. 2016). The Chinese paleosols offera global change baseline for comparison with long-term change in South Asian monsoonal paleosols,because they recordmiddle and lateMiocene spikesin precipitation and long-term drying also seen incomparable records fromNorthAmerica,Kenya,andAustralia (Retallack 2007a, 2013; Metzger and Re-tallack 2010).

A final pedogenic carbonate (Bk) metric is nodulesize, which has been calibrated in modern soils tothe age of surface soils (Retallack 2005). This dura-tion of soil formation is not the same as the geo-logical age of a paleosol but is useful for calculatingrates of sediment accumulation (Retallack 1997,2001). All three Bk metrics—depth, thickness, andnodule size—are here applied to the past 60 My ofChinese and South Asian paleosols as a novel andindependent record of Himalayan and Tibetan pa-leoclimate. Our study is the first to introduce pa-leosol Bkmetrics to understand the evolution of theSouth Asian monsoon. There are limits to furtherapplications of Bk metrics, because south Chineseand Southeast Asian Cenozoic paleosols are noncal-careous, as a result of humid to perhumid paleocli-mate (Nichols andUttamo 2005; Biasatti et al. 2012;Licht et al. 2014a). These simple fieldmetrics can beobtained rapidly and in volume, and we hope to en-courage similar studies in other regions of centralAsia and in time intervals missing from our study.Additional fieldwork is needed to extend our resultsto the thousands of red calcareous Cenozoic paleo-sols now known from Pakistan (Retallack 1991,1995), the Tibetan Plateau (Rowley andCurrie 2006;DeCelles et al. 2007), Mongolia (Sun and Windley2015), and Kyrgyzstan (Macauley et al. 2016).

Geological Background

Paleosols are well known in Himalayan forelandbasins, within the Subathu andDagshai Formations(Sangode et al. 2010; Singh et al. 2010; Srivastavaet al. 2013) and the Siwalik Group of India (Retal-lack 1991, 1995; Singh et al. 2011, 2012) and Nepal(Quade et al. 1989, 1995, 2003; Tanaka 1997). Thisstudy extends observations of paleosols to previouslystudied Cenozoic stratigraphic sections in Gansu,China, in intermontane basins north of the TibetanPlateau (Flynn et al. 1999; Fang et al. 2003; Li et al.

2 G . J . R E T A L L AC K E T A L .

2006). Paleosols were recognized by the trinity of (1)root traces (figs. 2F, 3G), (2) soil horizons (figs. 2C–2F,3B, 3C, 3G), and (3) soil structure (figs. 2D, 3C),as already documented for the Dagshai and Subathu

Formations of India (Singh et al. 2010; Srivastava et al.2013).Evidence of nonmarine paleoenvironments comes

from fossil land plants and terrestrial vertebrates in

Figure 1. Location of examined paleosols in Himachal Pradesh (India), Gansu (China), and other relevant paleosols insouthern Nepal, with enlargements of selected areas. Base map is from NASA (National Aeronautics and SpaceAdministration) Blue Marble winter view.

Journal of Geology 3S T RONG SOUTH A S I AN MON SOON B Y 2 0 MA

the Dagshai and Kasauli Formations (Mehra et al.1990;Arya et al. 2004; Bhatia andBhargava 2006) andall the sections described here from Gansu, China(Fang et al. 2003; Li et al. 2006; Qiu et al. 2013). Pol-len of palms and cells of freshwater algae in the basal

Dagshai Formation are evidence of ponded water andcoastal nonmarine habitats (Singh and Sarkar 1990).However, the Subathu Formation includes a mix ofmarine and nonmarine fossils and facies. Gray silt-stones of the Subathu Formation include fossil oyster

Figure 2. Field photographs of paleosols in Himachal Pradesh, India. A, Contact (at person) of Dagshai Formationsandstones (vertically dipping to left) and upper Subathu Formation (red beds) near Rehon (N30.844187 W77.088277).B, Kasauli Formation basal sandstone above red beds of Dagshai Formation dipping west, 3 km northeast of Kumar-hatti (N30.8885907 W77.069797). C, Chakki pedotype in middle Subathu Formation, 400 m east of Chakki-ka-More(N30.860267W77.004117).D, Kotla pedotype in upper Subathu Formation near Rehon (N30.841647W77.086637). E, Gahipedotype in lower Dagshai Formation, 1 km east of Chakki-ka-More (N30.860967 W77.0037287). F, Barog pedotype inupper Dagshai Formation, 3 km northeast of Kumar Hatti (N30.8885907 W77.069797).

4 G . J . R E T A L L AC K E T A L .

biostromes (Bhatia et al. 2013), large photosymbi-otic foraminifera (Bhatia et al. 2013), whales (Bajpaiand Gingerich 1998), and a distinctive assemblage ofsmall tropical sea shells (Mathur 1975). In contrast,calcareous red beds of the Subathu Formation in-clude freshwatercharophytes (BhatiaandBaghi1991)and snails (Mathur 1965) and terrestrial rodents attwodistinctstratigraphic levels (fig.4),earlyLutetian(middle Eocene) and late Ypresian (early Eocene;Srivastava and Kumar 1996; Gupta and Kumar 2015).Sites for red beds of the Subathu Formation visitedfor this work included paleosols, with root traces,soil horizons, and ped-cutan soil structure (fig. 2A,2C, 2D). These red beds do show subtle grain-sizegrading, but we do not agree with the interpretationby Bera et al. (2008) of Subathu red beds as marine“red calciturbidites.” Paleobathymetry of the marineSubathu Formation inferred from large photosym-biotic foraminifera and oyster biostromes was less

than 40 m (Singh 2012). Some Subathu fossil assem-blages have medium-sized (3–4 cm), smooth oysters(Flemingostrea flemingi) and little else. Other as-semblages of turritellid snails and carditid bivalvesare all less than 12mm long, and even the large fora-minifera are less than 15 mm in diameter; these dis-tinctive assemblagesmay record fresh to hypersalineconditions of coastal estuaries and salinas (Mathur1975). The basal Subathu Formation has coal seamswith deeply weathered tuffs (tonsteins) in Histosolpaleosols (Siddaiah and Kumar 2008) and bauxites ofOxisol paleosols (Singh et al. 2009). This is thefirst offour nonmarine levels in the largely shallow-marineSubathu Formation, below the three levels recog-nized in our composite section (fig. 4).Biostratigraphic ages for these various formations

are based primarily on large shallow-marine fora-minifera (Nummulites, Assilina) for the SubathuFormation (Bhatia et al. 2013), on palynology and

Figure 3. Field photographs of paleosols in Gansu, China. A, Late Miocene Yaodian Formation in gullies 2 km east ofYaodian (N34.617877 W105.903397). B, Trappo, Dam, and Mukpo pedotypes (top to bottom) at 45–47 m in Yaodiansection (N34.616127 W105.907777). C, Mukpo pedotype at 55–56 m in Yaodian section (N34.616327 W105.908217).D, Dokdo pedotypes at 70–80 m in Yaodian section (N34.616327 W105.908217). E, Striped orange beds of late MioceneDongxiang Formation above deeper red and white banded late Miocene Shangzhuang Formation and massive redZhongzhuang Formation, 8.0 km west of Dongxiang (N35.604437 E103.345687). F, Middle Miocene ZhongzhuangFormation, 4.0 km west of Dongxiang (N35.668887 E103.296247).G, Oligocene Thikpa pedotype with drab-haloed roottrace in Yehucheng Formation, near Diutinggou village (N36.2124497 E103.6213417). H, Sandy Xianshiuhe Formationabove and to right of clayey Yehucheng Formation, 2 km south of Diutinggou village (N36.214697 E103.612237). hozphorizon.

Journal of Geology 5S T RONG SOUTH A S I AN MON SOON B Y 2 0 MA

paleobotany for theDagshai andKasauli Formations(Mehra et al. 1990; Arya et al. 2004; Bhatia andBhargava 2006), and on mammalian biostratigraphyfor the Chinese formations (Flynn et al. 1999; Fanget al. 2003; Li et al. 2006). More precise age modelsfor the Chinese formations comes from paleomag-netic data (Fang et al. 2003; Li et al. 2006; Qiu et al.2013) updated to the timescale of Ogg (2012).

The age of the basal Dagshai Formation is con-troversial: 39.6 Ma, from paleomagnetic chron cor-relation (Sangode et al. 2010); 35.5 5 6.7 Ma, frompaleomagnetically determined paleolatitude (Naj-man et al. 1994); 32Ma, from fission-track dating ofdetrital zircons (Jain et al. 2009; Srivastava 2013);30Ma, fromU-Pb dating of detrital zircons (Najman2006; Ravikant et al. 2011); and 24.7 Ma, from 40Ar/39Ar dating of detritalmuscovite (Najman et al. 1997).An earlyMiocene age (20–23Ma) is likely from litho-logical correlation with the paleomagnetically datedDumri FormationofNepal (Ojha et al. 2009;DeCelleset al. 2014) and mammal faunas of the DharmsalaFormation of Kangra Valley, Himachal Pradesh (Ti-wari et al. 2006; Sehgal and Bhandari 2014), and theMurree Formation of Jammu and Kashmir (KumarandKad2003).Thedetritalmineral ages aremaxima,and the oldest estimate of 39.6 Ma comes from theuntenable assumption that the red upper SubathuFormation is transitional into the red lower DagshaiFormation (Singh and Lee 2007; Sangode et al. 2010;Bhatia et al. 2013). Regional mapping has revealed amajor geological disconformity, representedbyquartzsandstones, above the middle Eocene (early Lutetian,or at least 48 Ma) marine portion of the Subathu For-mation (Bera et al. 2008). A comparable large uncon-formity between marine shales and nonmarine redbeds has been found in Nepal (DeCelles et al. 2004,2014). A misleading appearance of transition in Hi-machal Pradesh comes frommiddle Eocene (early Lu-tetian) red beds with caliche nodules and terrestrialvertebrates (Singh and Lee 2007), such as red terres-trial mudstone lenses lower in this largely marineformation (Srivastava and Kumar 1996; Gupta andKumar 2015). What remains of the Dagshai Forma-tionmagnetostratigraphy of Sangode et al. (2010) is along normal above a short reversal, which we corre-late with chrons C6n (18.748–19.722 Ma) and C6r(19.722–20.04 Ma) of Ogg (2012). Unlike other cor-

Figure 4. Composite geological section of early Ceno-zoic sedimentary rocks of Himachal Pradesh, India (Suba-thu,Dagshai, andKasauli Formations) based on exposuresnear Subathu (N30.355657 E77.006757), Chakki-ka-More(N30.860867 E77.005147), and Kumarhatti (N30.871367E77.069867). Informal units A1–D18 are those of Singhet al. (2010) and Srivastava et al. (2013). Pedotype namesare to the left of the development box, based on the soil

maturity scale of Retallack (2001). Calcareousness as aguide to aridity is based on field reaction with diluteHCl, and hue as a guide towaterlogging is based on aMun-sell color chart. A color version of this figure is availableonline.

6 G . J . R E T A L L AC K E T A L .

relations of short normals with unreasonable sedi-ment accumulation rate, this correlation gives a sed-iment accumulation rate of 0.13 mm/y, comparablewith that of the basal Dumri Formation of Nepal(Ohja et al. 2009).Insecure dating of the Dagshai Formation remains

a problem, but the disconformity of over 20 My forour time series in Himachal Pradesh and Nepal maynot apply to all outcropsof red beds in theHimalayanforeland. Different measured sections of the DagshaiFormation in different fault blocks have different se-quences (Bera et al. 2010), and paleomagnetic and de-trital mineral dating studies in blocks other than theChakki-ka-More section may find some of the geo-logical time now consideredmissing.Paleolatitude of the Subathu region of northern

India was about 137N 45 My ago, but Indian north-ward movement slowed to transit 197–307N over thepast 35My, whereas north China andMongolia havebeen between 307N and 407N for the past 60 My(Najmanet al. 1994;Dupont-Nivet et al. 2010;Huanget al. 2015). Thus, the Subathu Formation formedsouth of 107N, well south of its current outcrop at317N, but the Dagshai Formations formed north of207N.Gansuwas at about at 367Nfor the past 60My.

Material and Methods

This work proceeded by reoccupying measured sec-tions in Himachal Pradesh, India (figs. 4, 5), andGansu, China (figs. 6–9), and logging all paleosolsencountered. Each newly discovered pedotype wasgraphically characterized, including Munsell colorand rock texture, and named after local villages inHimachal Pradesh or with simple descriptive termsin the Tibetan language (Bell 1989) for paleosols inGansu. Also gathered were field measurements offeatures of calcareous paleosols that have paleoen-vironmental significance (Retallack 2005): (1) depthto the calcareous nodular (Bk) horizon, (2) thicknessof paleosol with carbonate nodules, and (3) size ofnodules. These were all measured directly on theoutcropwithacloth tape,usuallyafter excavation todetermine the upper bounds of the paleosol. Depthand thickness of Bk ranged from 20 cm tomore than140 cm and nodule horizontal diameter from 0.1 to6 cm, with five exceptions up to 14 cm and startingto merge into a calcareous bench (K horizon). Ournew field observations are tabulated in the appen-dix, available online.Our new field data were supplemented with pre-

viously published geochemical data on Indian andNepali Cenozoic paleosols, including data on oxy-gen and carbon isotopic composition of pedogenic

carbonate nodules (Harrison et al. 1993;Quade et al.1995; Srivastava 2001; Ghosh et al. 2004; Sanyalet al. 2005; Singh and Lee 2007; Leier et al. 2009; Beraet al. 2010) and major-element chemical analysis ofnoncalcareous (!3 wt% CaO) parts of the paleosols(Najman and Garzanti 2000; Hossain et al. 2008;Siddaiah and Kumar 2008; Singh et al. 2009; Sri-vastava et al. 2013) and modern soils and sediments(Srivastava 2001; Singh 2010).

Paleosol Interpretation

Identification of the various pedotypes in tables 1and 2 used three separate soil classifications: (1) theupdated soil taxonomy of the United States Natu-ral Resources Conservation Service (Soil SurveyStaff 2014), (2) the international classification of theUnitedNations Educational, Scientific and CulturalOrganization (UNESCO; FAO 1977, 1978), and (3) atraditional classification of the Commonwealth Sci-entific and Industrial Research Organization of Aus-tralia (Stace et al. 1968). The FAO (Food and Agri-culture Organization)-UNESCO classification (FAO1977, 1978) is particularly useful because it is appliedconsistently in classifying all land surfaces of theEarth intomapunits definedby a particular code. Forexample, map unit Bk40-2a (Je, Jc, G) near Muza-farpur, India (FAO 1977), consists mainly of CalcicCambisols (Bk), with subordinate Eutric Fluvisols,Calcaric Fluvisols, and Gleysols. Individual geolog-ical formations and parts of formations are assignedcomparable codes in table 3, which lists the mostsimilar modern map units. This is thus a guide tocomparable modern soil-forming environments, orpedoscapes, giving a paleoenvironmental context.Interpretations of likely paleoenvironmental fac-

tors for each pedotype (tables 1, 2) are based on pre-viously published studies of these particular paleo-sols (Retallack1991, 1995;Quade et al. 1995;Tanaka1997; B. P. Singh et al. 2010; S. Singh et al. 2012;Srivastava et al. 2013) or other guides to paleosol in-terpretation (Retallack 1997, 2001). In the EoceneSubathu Formation, mangal paleosols colonizingoyster biostromes are revealed by both carbonaceousroot traces in oyster coquinas and oysters (Fleming-ostrea flemingi) with umbonal deformation (Mathur1975), as if growing around narrow mangrove proproots (Plaziat 1970). Coastal floodplain paleosols ofthe Subathu, Dagshai, and Kasauli Formations havedrab-haloed root traces (fig. 2C), as if they were orig-inallywell drained but subsided into a shallowwatertable after burial (Retallack 1991).Quantitative paleoclimatic estimates can be de-

rived from a previously published global database of

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Figure 5. Paleosols of Himachal Pradesh, India. A, Lower Dagshai Formation, road cut 8 km east of Kumarhatti(N30.849147 E77.08757). B, Lower Kasauli Formation, road cut 4 km east of Kumarhatti (N30.864347 E77.074447). C,Lower Dagshai Formation, road cut 2 km east of Chakki-ka-More (N30.860967 E77.0037287). D, Lower DagshaiFormation, road cut 1.5 km east of Rehon (N30.844187 E77.088277). E, Lower Kasauli Formation, road cut 2.0 kmnortheast of Kumarhatti (N30.889757 N77.066927). F, Upper Subathu Formation, road cut 1.0 km east of Rehon(N30.849147 N77.087507). G, Middle Subathu Formation, road cut 2 km east of Chakki-ka-More (N30.860267E77.004117). H, Lower Subathu Formation north of Kuthar River, 400 m east of bridge, 1.6 km southwest of Subathu(N30.964817 N76.97967). A color version of this figure is available online.

674 soils of postglacial age in unconsolidated sedi-ments of low-lying terrane under grasslands andshrublands (Retallack 2005), which found a relation-shipbetweenmeanannualprecipitation(P inmm)and

depth to calcareous nodules (D in cm), given by thefollowing formula, with standard error of 5147 mmand a coefficient of determination (r2p 0.52):

P p 137:241 6:45D2 0:013D2: ð1Þ

The standard error accommodates data from allclimates for universal applicability to paleosols, in-cluding nine soils from India and Pakistan and foursoils from northwest China. This relationship hasbeen questioned by Royer (1999), who used a data-base of bedrock soils inappropriate to sedimentarysettings (Retallack 2000). Not only topographic set-ting but parentmaterials, vegetation, and formationdurations must be constrained for this and otherclimofunctions, as stressed when first proposed byJenny and Leonard (1935). The climofunction wasnotbasedonsoils thatwerewaterlogged (unoxidizedwith siderite or pyrite) or developed on bedrockslopes (with metamorphic or igneous rock below),under rain forests (thick kaolinitic and noncalcar-eous), or for longer than about 10,000 y (so with car-bonatebenchesorKhorizons;Retallack1997,2001).Thus, the climofunction was not applied to paleo-sols with any of these distinctive features but waslimited to oxidized and freely rooted and burrowedpaleosols of alluvial sediments, illitic to smectiticcomposition, and small nodules.Also from the same study and database (Retallack

2005), the relationship between the mean annualrange of precipitation (M as mean precipitation inmm of the wettest minus the driest month) and thethickness of soil between the highest and lowestnodules (T in cm) is given by equation (2), with astandard error of 522 mm and a coefficient of de-termination (r2 p 0.58). The nine soils from Indiaand Pakistan in the database for this relationshipare the most extreme for seasonality of precipita-tion, a key component ofmonsoon intensity (Huberand Goldner 2012; Wang and Chen 2014), and thusdefine the relationship:

M p 0:79T 1 13:71: ð2Þ

Both proxies were corrected for compaction ofpaleosols due to burial, using equation (3) for Ari-disols (Sheldon and Retallack 2001), giving depth orthickness of horizon with nodules in the originalsoil (Ds in cm) from depth in the current paleosol(Dp in cm) for a particular depth of burial (K in km).Depth of burial was taken from local stratigraphiccompilations (Quade et al. 1995; Tanaka 1997; Flynnet al. 1999; Fang et al. 2003; Li et al. 2006; B. P. Singhet al. 2010; S. Singh 2012; Srivastava et al. 2013), andthe Aridisol function was used because a calcic hori-

Figure 6. Composite sections in Gansu, China: A, gullysoutheast of Yaodian (Li et al. 2006); B, road cuts aroundDongxiang (Fang et al. 2003); C, road cuts near Diu-tinggou (Qiu et al. 2013). Pedotype names are to the leftof the development box, based on the soil maturity scaleof Retallack (2001). Calcareousness as a guide to aridityis based on field reaction with dilute HCl, and hue as aguide to waterlogging is based on a Munsell color chart.Chron ages are from Ogg et al. (2012). A color version ofthis figure is available online.

Journal of Geology 9S T RONG SOUTH A S I AN MON SOON B Y 2 0 MA

zon at a depth of less than 1m is part of the definitionof Aridisols (Soil Survey Staff 2014):

Ds pDp

20:62=½(0:38=e0:17K)2 1� : ð3Þ

Another relationship documented in the samestudy of modern soils (Retallack 2005) is that be-tween duration of soil formation (A in ky) and thediameter of calcareous nodules (S in cm), given byequation (4), with a standard error of 51.8 ky and acoefficient of determination (r2 p 0.57). This rela-tionship is based on only 10 radiocarbon-dated soilsfrom New Mexico (Gile et al. 1981), still the mostcomprehensive available database. Compaction cor-rection is not applied to horizontal diameter of nod-ules or logs, which cannot spread because of laterallithostatic support (Retallack 1994):

A p 3:92S0:34: ð4Þ

The complex petrography of South Asian nodulescomparable with modern monsoonal soils has beendescribed elsewhere (Retallack 1991), and here weadd observations of rooting depth and the depth,thickness, andnodule sizeof thecalcichorizon (fig.2)and compare themwith those of other paleosols pre-viously described from India andNepal (Quade et al.1989, 1995; Retallack 1991; Sangode et al. 2010; Sri-vastava et al. 2013). Depth to calcic horizon andthickness of paleosols with carbonate could both becompromised by erosion of profiles before burial, sothat our data canbe regarded asminimal estimates ofprecipitation and seasonality. With the exception ofthe basal Dagshai Formation disconformity (fig. 2A,2D), contacts above the paleosols are nonerosive andpreserve large root traces (figs. 2C, 2E, 2F, 3). Apartfrom the sub-Dagshai unconformity, the sectionsexamined are thick and continuous, and other stud-ies of sedimentary facies and sedimentation ratesconfirm the aggradational nature of these unusually

Figure 7. Paleosols of Diutinggou, Gansu, China. A, Lower Xianshiuhe Formation, road cut 1 km west of Diutinggou(N36.215727 E103.610677). B, C, E, Upper Yehucheng Formation, successive paleosols in badland north of road, 0.8 kmwest of Diutinggou (N36.214697 E103.612237).D, F,G, Middle Yehucheng Formation in badland south of road, 0.25 kmwest of Diutinggou (N36.212927 E103.618757). H–J, Lower Yehucheng Formation, road cut immediately south ofDiutinggou (N36.2124497 E103.6213417). A color version of this figure is available online.

10 G . J . R E T A L L A CK E T A L .

thick alluvial sequences (Behrensmeyer et al. 1995;Flynn et al. 1999; Fang et al. 2003; Li et al. 2006;Ojhaet al. 2009; Srivastava et al. 2013).Stable-isotope data are also available on caliche

nodules of Bk horizons of Indian and Nepali paleo-sols, including 27 of the paleosols in this study(Leier et al. 2009; Bera et al. 2010). This reveals astriking change in vegetation from C3 woodland toC4 grassland at about 5–7 Ma (Quade et al. 1989).Vegetation change to sod grasslands is also sup-ported by field observations of crumb peds of Mol-lisol in paleosols at about 5–7 Ma in Pakistan andNepal (Quade et al. 1995; Retallack 2013). Using

similar soil-structural criteria, Mollisols appear at7.5 Ma in the Yaodian section and at 7 Ma in theDongxiang section (fig. 6). Following Fox and Koch(2003), absence of C4 grasses is indicated by valueslower than 210.1‰ d13C, 50% C4 grasses by 23.1‰d13C, and 100% C4 by 13.9‰ d13C. Oxygen isotopiccompositions of pedogenic nodules in the Himala-yan foreland are more difficult to interpret, becausethey have amix of source, temperature, rainout, andgeographic influences. Summer monsoon rains areintense and warm and have lower d18O than theirocean source (Licht et al. 2014b), but then lowertemperature and lower available moisture farther

Figure 8. Paleosols near Dongxiang, Gansu, China. A, Lower Hewangjia Formation, road cut on farm road belowridge, 6.4 km southwest of Dongxiang (N35.613017 E103.35027). B, F, Upper Hewangjia Formation, in badlands belowridge, 8.0 kmwest of Dongxiang (N35.668887 E103.296247).C–E, Liushu Formation, road cuts on farm road below ridge,6.3 km southwest of Dongxiang (N35.614097 E103.349927). G, Upper Shangzhuang Formation, road cut 2.8 kmsouthwest of Dongxiang (N35.648907 E103.362817).H, Middle Zhongzhuang Formation, road cut 3.2 km southwest ofDongxiang (N35.652117 E103.352977). I, Middle Tala Formation, road cut 7.7 km west of Dongxiang (N35.658477E103.300537). A color version of this figure is available online.

Journal of Geology 11S T RONG SOUTH A S I AN MON SOON B Y 2 0 MA

inland also lower the d18O of rain and snow (Araguás-Araguás et al. 1998).

Publishedmajor-element analyses of the paleosolsalso provide evidence of degree of chemical weather-ing by hydrolysis, which depletes cation bases (Ca21,Mg21, Na1, and K1) at the expense of alumina inclays. The weathering proxy chosen for this studyand by Clift et al. (2008) is the chemical index of al-teration (I in %) of Nesbitt and Young (1982), cal-culated frommolar proportions (m) of alumina, lime,potash, and soda according to equation (5). The limeis noncarbonate lime, so samples with 13% CaOwere excluded:

I p100(mAl2O3)

mAl2O3 1mCaO1mNa2O1mK2O: ð5Þ

The chemical index of alteration of deeply weath-ered tropical soils and sediments is greater than 80,but few weathered periglacial soils and sedimentshave values of less than 60 (Nesbitt and Young 1982;

Passchier and Krissek 2008; Bahlburg and Dobr-zinski 2011).

Paleoclimatic Record of Asian Monsoon

Our time series for paleoclimate of Asia is incom-plete (fig. 10) but offers evidence for a surprisingantiquity of highly seasonal precipitation charac-teristic of the modern South Asian monsoon. Theoldest known thick-calcic paleosol with the hall-marks of modernmonsoonal soils is in theMiocene(20 Ma) part of the Dagshai Formation of HimachalPradesh (figs. 2E, 6C), where the disconformablyunderlying Eocene Subathu Formation has thin-calcic paleosols like soils formed in less seasonalclimates (figs. 2D, 5F). Thin-calcic paleosols likethose of Eocene India and inGansu today are seen inChina over the past 34 My. Both mean annual pre-cipitation from depth to calcic horizon and meanannual range of precipitation from spread of car-bonate can be calculated with equations (1) and (2)

Figure 9. Paleosols in the Yaodian Formation (A–G) and an unnamed formation (H) in Yaodian gully, Gansu, China,observed in measured sections at 60 m (A; N34.616327 E105.908217), 55 m (B; N34.616327 E105.908217), 45 m (C;N34.616127 E105.907777), 90m (D; N34.616177 E105.909277), 116m (E; N34.616097 E105.910487), 161m (F; N34.616097E105.910487), 220 m (G; N34.616157 E105.911277), and 310 m (H; N34.6155497 E105.915897). A color version of thisfigure is available online.

12 G . J . R E T A L L A CK E T A L .

to quantify these observations (fig. 10A, 10B). Pa-leosols of the 20 Ma part of the Dagshai Formationrecord a modern range of monsoon seasonality inIndia,whereas the less seasonal climate of the 48MaSubathu Formation is comparable to that of north-west China over the past 34 My (fig. 10B). The in-ferred difference between dry- and wet-month pre-cipitation from Bk metrics is less than half thatdemonstrated inmodern environments, because themostextrememodernseasonality is insoils sohumidthat they lack carbonate, such as Neogene paleosolsof the eastern Himalaya and Southeast Asia (Nich-ols andUttamo 2005; Biasatti et al. 2012; Licht et al.2014a). Our incomplete paleosol time series is thuscompatible with the Eocene (40 Ma) onset of mon-soonal seasonality inferred from oxygen isotopic var-iation in growth rings of fossil snails and mammals(Licht et al. 2014b) and fossil wood anatomy (Lichtet al. 2015) and with modeling studies suggestingSouth Asian monsoons as early as 45–35 Ma (Huberand Goldner 2012) or 30 Ma (Ramstein et al. 1997).Our evidence for nonmonsoonal seasonality duringthe early Eocene is supported by fossil floras of thatage in southern China (Spicer et al. 2016).Shallow calcic horizons of paleosols in the lower

Dagshai Formation are evidence of semiarid paleo-climate like that of the Oligocene-Holocene Gansuand the Eocene Subathu Formation, but subhumidpaleoclimate is apparent from deep calcic paleosolsof the upper Dagshai and lower Kasauli Formations

(fig. 10A). Thus, the increase in mean annual precip-itation revealed by increased depth to Bk horizonspostdated the increase in seasonality of precipitationfrom increased thickness of Bk horizons. Synorogenicdeposition of the Dagshai and Kasauli Formations(Bera et al. 2008; Srivastava et al. 2013) by 20 Mathus coincided with an increase in paleoprecipita-tion in the north Indian foreland. These precipita-tion increases explain enhanced chemical weather-ing evident from the chemical index of alteration ofmarine sedimentary rocks of the Arabian Sea (Cliftet al. 2008).Secular variation in precipitation after 20 Ma

(fig. 10A) also tracks sedimentary weathering indi-ces (Sun and Windley 2015), without a local changein facies of silty red beds (Quade et al. 1995; Tanaka1997). Increases in mean annual precipitation atabout 5 and 10 Ma (fig. 10B) correspond to precipi-tation spikes in western North America (Retallack2007a), Kenya (Retallack 2007b), and South Austra-lia (Metzger and Retallack 2010). The 10 Ma spikecoincides with diatom blooms in the Indian Ocean(Gupta et al. 2004) and the 5Ma spike with increaseddust influx inmarine cores (An et al. 2001). These areboth times of higher atmospheric CO2, from sto-matal index of fossil oak (Kürschner et al. 1996) andginkgo leaves (Retallack 2013), and they support thegeneral notion that high CO2 encourages higher wa-ter vapor and a more active hydrological cycle (Re-tallack 2009). These spikes correspond to global

Table 1. Interpretation of Pedotypes, Himachal Pradesh, India

Pedotype Paleoclimate Ecosystem Paleotopography Parent materialTime for

formation (y)

Barog Subhumid mon-soonal tropical

Monsoonwoodland

Well-drainedfloodplain

Quartzofeldspathicsilt

5000–6000

Chakki Not diagnostic Early successionalpole woodland

Well-drainedfloodplain

Quartzofeldspathicsilt

200–500

Dharampur Not diagnostic Mangrovewoodland

Intertidal flats andoyster reefs

Oysters and clay 10–100

Gahi Semiarid mon-soonal tropical

Monsoon woodland Well-drainedfloodplain

Quartzofeldspathicsilt

5000–6000

Rehon Not diagnostic Swamp forest Seasonally wetfloodplain

Quartzofeldspathicsilt

2000–3000

Kumar Not diagnostic Early successionalswamp woodland

Seasonally wetswampyfloodplain

Quartzofeldspathicsilt

500–1000

Koshalia Not diagnostic Early successionalshrubland

Well-drainedstreamsidelevee

Quartzofeldspathicsilt

10–100

Koti Not diagnostic Swamp forest Seasonally wetfloodplain

Quartzofeldspathicsilt

500–1000

Kotla Semiaridnonmonsoonaltropical

Dry tropicalwoodland

Well-drainedfloodplain

Quartzofeldspathicsilt

2000–3000

Solan Not diagnostic Swamp forest Seasonally wetdry floodplain

Pisolitic and quartz-feldspathic silt

2000–3000

Journal of Geology 13S T RONG SOUTH A S I AN MON SOON B Y 2 0 MA

Tab

le2.

Interpretation

ofPed

otyp

es,Gan

su,China

Ped

otyp

ePaleo

clim

ate

Eco

system

Paleo

topo

grap

hy

Paren

tmaterial

Tim

efor

form

ation(y)

Chhem

aNot

diag

nos

tic

Early

succ

ession

alsh

rublan

dWell-drained

alluvial

leve

eQuartzofelds

pathic

silt

10–10

0Dam

Subh

umid

Tallgrasslan

dwithscattered

tree

sWell-drained

floo

dplain

Quartzofelds

pathic

silt

andclay

100–

2000

Doh

ruSe

miaridtempe

rate

Desertsh

rublan

dWell-drained

alluvial

fan

Quartzofelds

pathic

grav

el10

0–20

00Dok

doSe

miaridtempe

rate

Desertsh

rublan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

2000

–40

00Dre

Semiaridtempe

rate

Dry

woo

dlan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

2000

–30

00Dru

Not

diag

nos

tic

Early

succ

ession

alsh

rublan

dWell-drained

stream

side

leve

eQuartzofelds

pathic

sandan

dsilt

10–10

0Dza

no

Not

diag

nos

tic

Early

succ

ession

alsh

rublan

dWell-drained

stream

side

leve

eQuartzofelds

pathic

silt

100–

2000

Gya

Subh

umid

tempe

rate

Tallgrasslan

dWell-drained

floo

dplain

Quartzofelds

pathic

calcareo

us

loess

2000

–30

00

Gyu

paHyp

erarid

toarid

Halop

hytic

shru

blan

dDay

salt

pan(playa

)Quartzofelds

pathic

silt

100–

500

Jangk

hu

Not

diag

nos

tic

Carr(alkalinewetland

woo

dlan

d)Se

ason

ally

dryfloo

dplain

wetland

Lac

ustrinemarlan

dsilt

10–10

0

Khung

Subtropica

lEarly

succ

ession

alsh

rublan

dwithterm

itemou

nds

Well-drained

alluvial

leve

eQuartzofelds

pathic

sand

10–10

0

Kupa

Subh

umid

tempe

rate

Dry

woo

dlan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

2000

–40

00Kupk

yaNot

diag

nos

tic

Grassland

Well-drained

floo

dplain

Quartzofelds

pathic

calcareo

us

loess

20,000

–50

,000

Map

oSu

bhumid

tempe

rate

Tallwoo

dlan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

2000

–50

00Mukpo

Semiaridtempe

rate

Grassland

Well-drained

floo

dplain

Quartzofelds

pathic

calcareo

us

loess

2000

–30

00

Ngo

pkya

Semiaridtempe

rate

Fen(alkalinewetland

grasslan

d)Se

ason

ally

waterlogg

edfloo

dplain

Quartzofelds

pathic

calcareo

us

loess

2000

–30

00

Pan

gNot

diag

nos

tic

Grassypo

lewoo

dlan

dSe

ason

ally

waterlogg

edfloo

dplain

Quartzofelds

pathic

calcareo

us

loess

1000

–20

00

Sakar

Semiaridtempe

rate

Desertsh

rublan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

50,000

–10

0,00

0Se

poSe

miaridtempe

rate

Desertsh

rublan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

2000

–40

00Sh

eAridtempe

rate

Desertsh

rublan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

1000

–20

00Sh

oko

Subh

umid

tempe

rate

Dry

woo

dlan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

2000

–40

00Sh

ompa

Not

diag

nos

tic

Succ

ession

alpo

lewoo

dlan

dWell-drained

stream

side

leve

eQuartzofelds

pathic

silt

10–10

0Thikpa

Semiaridtempe

rate

Desertsh

rublan

dWell-drained

alluvial

fan

Quartzofeldspa

thic

sand

1000

–20

00Thukpo

Subh

umid

tempe

rate

Dry

woo

dlan

dWell-drained

floo

dplain

Quartzofelds

pathic

silt

2000

–50

00Trapp

oNot

diag

nos

tic

Early

succ

ession

alsh

rublan

dStream

side

leve

eClay

10–10

0Tsa

Not

diag

nos

tic

Carr(alkalinewetland

woo

dlan

d)Se

ason

ally

dryfloo

dplain

wetland

Quartzofelds

pathic

silt

2000

–30

00

Tsawa

Not

diag

nos

tic

Succ

ession

alpo

lewoo

dlan

dInterm

ittentlydrained

stream

side

leve

eQuartzofelds

pathic

siltan

dsand

100–

1000

14

Tab

le3.

Interpretation

ofPed

oscape

sin

Indiaan

dChina

Form

ation

Cou

ntry

Age

(Ma)

FAO

(197

7,19

78)

code

forform

ation

Mos

tsimilar

FAO

(197

7,19

78)

mod

ernco

deMod

ernen

vironmen

t

Upp

erSiwalik

Group

India

2–4

Bk(C

k,Jc,Gc)

Bk40

-2a(Je,

Jc,G);Muza

farpur,India

Mon

soon

alsu

btropica

lclim

ate

Upp

erHew

angjia

Form

ation

China

5–4

Ck

Ck1-3b

;Balpy

kBi,Kaz

akhstan

Warm

continen

talde

sert

Low

erHew

angjia

Form

ation

China

6–5

Ck

Ck1-3b

;nea

rBalpy

kBi,Kaz

akhstan

Warm

continen

tal

desert

MiddleSiwalik

Group

Nep

al8–

6Lc(Bk,Jc)

Lc7

3-2b

c(Be,

I,Je,G);

Mah

iyan

ganay

a,SriLan

ka

Hot,h

umid

mon

soon

alclim

ate

Upp

erYao

dian

Form

ation

China

8–6

Bk(Be,

Gc)

Bk44

-3a(Xk,Gc,

Je);Hen

gshan

,Sh

aanxi,China

Semiw

arm

continen

talde

sert

Liush

uFo

rmation

China

8–6

Ck(G

c)Ck3-3a

;Step

nya

k,Kaz

akhstan

Warm

continen

talde

sert

MiddleYao

dian

Form

ation

China

10–8

Xk(Be,

Rc,

Gc)

Xk37

-2ab

(Zo);Baiyin,Gan

su,China

Semiw

arm

continen

talde

sert

Upp

erDon

gxiangFo

rmation

China

11–8

Bk(G

e,Xk)

Bk36

-2b(Lo,

Xk,Rc);Dingy

uan

zhen

,Gan

suChina

Warm

continen

talde

sert

Low

erSiwalik

Group

Nep

al12

–10

Lc(Bk,Jc)

Lc7

3-2b

c(Be,

I,Je,G);

Mah

iyan

ganay

a,SriLan

ka

Hot,h

umid

mon

soon

alclim

ate

Low

erYao

dian

Form

ation

China

12–10

Xk(Rc,

Gc)

Xk37

-2ab

(Zo);Baiyin,Gan

su,China

Semiw

arm

continen

talde

sert

Low

erDon

gxiangFo

rmation

China

13–11

Xk(Bk,Gc)

Xk37

-2ab

(Zo);Baiyin,Gan

su,China

Semiw

arm

continen

talde

sert

Shan

gzhuan

gFo

rmation

China

14–15

Xk(Be,

Je)

Xk37

-2ab

(Zo);Baiyin,Gan

su,China

Semiw

arm

continen

talde

sert

Zhon

gzhuan

gFo

rmation

China

18–15

Lc(Be,

Je)

Lc1

01-2a(Lo,

I,Je);Tan

chen

g,Sh

ando

ng,

China

Mon

soon

warm

continen

tal

step

peLow

erKasau

liFo

rmation

India

21–22

Lc(Be,

Je,Ge)

Lc7

3-2b

c(Be,

I,Je,G);

Mah

iyan

ganay

a,SriLan

ka

Hot,h

umid

mon

soon

alclim

ate

TalaFo

rmation

China

22–21

Xk(Be,Je)

Xk37

-2ab

(Zo);Baiyin,Gan

su,China

Semiw

arm

continen

talde

sert

Upp

erDag

shai

Form

ation

India

25–24

Lc(Be,

Je,Gh)

Lc7

3-2b

c(Be,

I,Je,G);

Mah

iyan

ganay

a,SriLan

ka

Hot,h

umid

mon

soon

alclim

ate

Low

erDag

shai

Form

ation

India

25–26

Xk(Be,

Je,Vc)

Xk19

-2a(Zo);nea

rFe

rozepu

r,India

Hot

mon

soon

alsemitropical

clim

ate

Grave

lsin

base

ofYao

dian

gully

China

26–24

Xk

Xk1-2a

;Weiyu

an,Gan

su,China

Semiw

arm

continen

talde

sert

Low

erXiansh

uiheFo

rmation

China

29–28

Bk(G

e)Bk36

-2b(Lo,

Xk,Rc);Dingy

uan

zhen

,Gan

suChina

Warm

continen

talde

sert

Upp

erYeh

uch

engFo

rmation

China

30–29

Lk(Xk,Bk,Re,

Je)

Lc1

01-2a(Lo,

I,Je);Tan

chen

g,Sh

ando

ng,

China

Mon

soon

warm

continen

tal

step

peLow

erYeh

uch

engFo

rmation

China

33–31

Yy(Zo,

Fe)

Yy1

4-a(Zo);Pai-tun-tzu

,Gan

su,

China

Warm

continen

talde

sert

Upp

erSu

bath

uFo

rmation

India

49–48

Xk(Be,

Je,Jt)

Xk4-1b

;Belek

,Turk

men

istan

Con

tinen

talsemiaridMed

iter-

ranea

nclim

ate

Low

erSu

bath

uFo

rmation

India

52–49

Xk(Be,

Je)

Xk4-1b

;Belek

,Turk

men

istan

Con

tinen

talsemiaridMed

iter-

ranea

nclim

ate

15

Figure 10. A–C, Temporal changes in mean annual precipitation (A), seasonality of precipitation (B), and duration ofsoil formation (C) inferred from calcic horizon metrics of paleosols from Himachal Pradesh (India), compared withthose of Gansu (China) over the past 60 My. D–F, Pedogenic stable-isotope data (D, E) are from Harrison et al. (1993),Quade et al. (1995), Srivastava (2001), Ghosh et al. (2004), Sanyal et al. (2005), Singh and Lee (2007), Leier et al. (2009),and Bera et al. (2010), and chemical index of alteration (F) was calculated from data of Najman and Garzanti (2000),Hossain et al. (2008), Siddaiah and Kumar (2008), Singh et al. (2009), and Srivastava et al. (2013). Dashed gray linesrepresent epoch boundaries. Error bars are all 1 standard deviation from the transfer functions. G–I, Paleoclimatic andsoil duration interpretations were based on field measurements of calcic horizons of paleosols.

atmosphericchanges rather thanregionalorographicmonsoonal circulation.Paleosol durations are inversely related to sedi-

mentation rate, and truncations in paleosol-durationtrends reveal sequence boundaries (Retallack 1998;Mack andMadoff 2005; Atchley et al. 2013). Our ap-plication of the chronofunction of equation (4) revealsa high sedimentation rate throughout the sequencein India but rather slower sedimentation in Gansu,China (fig. 10C). These inferences are supported byrates of sedimentation from age models based onmagnetostratigraphy: 0.13 mm/y for the DagshaiFormation of India but 0.028 mm/y for Diutinggou,0.031 mm/y for Dongxiang, and 0.021 mm/y forYaodian. Furthermore, there was rapidly acceleratedsedimentation during deposition of the DagshaiFormation and then the Kasauli Formation. As pa-leosol duration declined in the Himalayan foreland,the chemical index of alteration in the paleosolsthere also declined to levels that reflect the glaci-atedmountainous source regions of these sediments(fig. 10F).

Himalaya or Tibetan Plateau Cause of Monsoon?

The Himalayan Range is such a striking barrier thatonewould think the summermonsoonwas enhancedby that great obstacle (Boos and Kuang 2010; Molnaret al. 2010). However, a wider view of South Asiancirculation emphasizes the importance of theTibetanPlateau for inducing monsoonal seasonality (Wuet al. 2012). Both the elevation of theTibetanPlateauand its displacement of Paratethyan seas have beenshown to inducemonsoonal circulation in computermodels (Fluteau et al. 1999; Roe et al. 2016).Our new data fail to address the relative role of the

Tibetan Plateau and Himalaya in generating highlyseasonal precipitation, because both were very highby the 20 Ma advent of monsoonal paleosols in theDagshai Formation, and there is currently no Indianpaleosolrecordbetween20and48Ma.Thesedimentsexamined here are now uplifted, folded, and faultedwithin the Lesser Himalaya but lack the thick syn-orogenic conglomerates of the upper Siwalik Groupas evidence of mountains close at hand (Bera et al.2010; Srivastava et al. 2013), as expected from pal-inspastic restoration of depocenters (DeCelles et al.2014). The sedimentary style and paleosols studiedhere are comparable to those of the Indo-Gangeticplains rather than the proximal Himalayan alluvialfans (Retallack 1991, 1995). The basal Dagshai For-mation has fossil algae and palm pollen of a coastalsetting (Singh and Sarkar 1990) and formed at a timewhen the Paratethyan Ocean was not completely

expelled from the Katawaz Basin of Pakistan to thewest (Patnaik 2016).Our evidence for monsoonal seasonality of pre-

cipitation by 20 Ma is close in time to radiometricand structural evidence for increased exhumationof the High Himalaya beginning at 23 Ma (Thiedeet al. 2004; Clift et al. 2008). Oxygen isotopic com-position of hornblende and biotite has been takenas evidence that Mount Everest and the Himalayanmain range were at elevations of 5 km by 17 Ma(Gébelin et al. 2013). However, the Tibetan Plateauis older than 26Ma, from comparable evidencemar-shaled by DeCelles et al. (2007), and coeval withOligocene desertification of central Asia (Sun andWindley 2015; Zheng et al. 2015). Isotopic evidencefrom paleosols on the Tibetan Plateau has been usedto infer 4 km of elevation by 35 Ma (Harris 2006;Rowley and Currie 2006; Currie et al. 2016; Ingallset al. 2016, 2017), and fossil leaf assemblages areevidence of 5 km of elevation since 15 Ma (Spiceret al. 2003). A recent review (Tada et al. 2016) out-lines three regional pulses of the Tibetan Plateau up-lift: starting in the south and central region at 40–35Ma, in thenorth at 25–20Ma, and in thenortheastand east at 15–10 Ma. The primacy of the TibetanPlateau for Indianmonsoonal climate could be testedwith discovery of calcareous paleosols in the Indianforeland dated to 25Ma or older.The Mongolian Plateau also induces East Asian

monsoonal circulation (Sha et al. 2015), but mon-soons are not, and have not been, intense in Gansu(fig. 10). The East Asian monsoon has been attrib-uted to seasonal migration of the intertropical con-vergence zone and moist static air over warm seas(Roe et al. 2016; Spicer et al. 2016).

Global Cooling and Drying byHimalayan-Tibetan Uplift?

Our new paleoclimatic data are relevant to the hy-pothesis of Raymo and Ruddiman (1992) that globalclimatic cooling and lowered precipitation over thepast 40 My were due to mountain uplift in generaland the Himalaya in particular. TheMiocene rise ofhot, humid monsoonal Indian climate thus contra-dicted the global trend represented here by Chinesedata (fig. 10). Using evidence from strontium isoto-pic ratios in marine carbonates, Raymo and Ruddi-man (1992) envisaged chemicalweathering enhancedby mountain uplift consuming CO2 from the atmo-sphere. Anderson et al. (2007) andMaher and Cham-berlain (2014) also consider physical and chemicalweatheringenhancedbyuplift.Thesehypothesesarecontradicted by evidence presented here of calcare-ous paleosols with minimal chemical weathering in

Journal of Geology 17S T RONG SOUTH A S I AN MON SOON B Y 2 0 MA

Himalayan and Tibetan flanking basins (figs. 4–9).Our data show semiarid to subhumid paleosols inIndia and China over the past 52 My (fig. 10A), incontrast to deeply weathered tonsteins (Siddaiah andKumar 2008) and bauxites (Singh et al. 2009) beforethat time (fig. 10F). Similarly, paleosolsasoldas35Maon the Tibetan plateau are calcareous and weaklyweathered in cold, arid paleoclimate (Rowley andCurrie 2006; Saylor et al. 2009). Onset ofmonsoonalseasonality, from Bk thickness (fig. 10B), was notinitially accompanied by an increase inmean annualprecipitation, from Bk depth in India, although pre-cipitation increase came soon afterward (fig. 10A).There has also been a decline in paleosol maturitysince20MainIndiabutnot incentralnorthernChina(fig. 10C), as rates of sedimentation increased. Thismatches the declining chemical index of alterationin paleosols of the Himalayan foreland (fig. 10F),which sinks below tropical values of 80 by 58 Ma,then below temperate values of 70 by 48 Ma. Theselowvaluesmayreflectanorigininglacier-fedstreamsfrom the High Himalaya, only partly offset by flood-plain weathering (Lupker et al. 2012). These ob-servations support the idea of increased physicalweathering but decreased chemical weatheringwithHimalayan and Tibetan uplift: it is chemical, notphysical, weathering that consumes atmosphericCO2. Conflation of physical and chemical weather-ing is the fatal flaw of arguments of Raymo andRuddiman (1992), Anderson et al. (2007), andMaherand Chamberlain (2014).

Mountain uplift inhibits chemical weathering inseveral ways, primarily by inhibiting plants, and isa force for global warming, not cooling. The stron-tium isotopic shift observed in Himalayan streamsis due not to hydrolyticweathering but to carbonatedissolution, which is carbon neutral for the atmo-sphere (Jacobson et al. 2002;Quade et al. 2003).Mod-ern soils and sediments of the Himalayan forelandare also barely weathered (Srivastava 2001; Singh2010), like the paleosols. Furthermore, metamor-phic decarbonation reactions with tectonic uplift ofthe Himalaya are degassed through widespread hotsprings; for example, the 32,000-km2 Narayani Ba-sin of the central Himalaya is degassing 11.3# 1010

mol/y CO2, which is four times the consumption ofCO2 by chemical weathering in that basin (Evanset al. 2008).

Something other thanHimalayan uplift promotedthe deeperweathering revealed bymarine strontiumisotopic values over the past 34My, and it is unlikelyto be completion of the circum-Antarctic current, asfirst Australia and then South America drifted awayfrom Antarctica with seafloor spreading (Kennett1982). Thermal isolation of Antarctica had an effect

comparable to that of Tibetan glaciation in curtail-ing chemical weathering in Antarctica (Retallacket al. 2001). The albedo effect of Antarctic ice wasnot sufficient to cool the world without help fromdeclining CO2 in the atmosphere, which Antarcticaworked against as its carbon-sequestering ecosys-tems were replaced by ice (DeConto and Pollard2003).

A more likely explanation for Cenozoic globalcooling is coevolution of grasses and grazers, asexplained in detail elsewhere (Retallack 2013). Or-ganismswithin coevolutionary trajectories are con-cerned less with their physical environment thanwith their biological environment, and this givescoevolution a potential for changing physical envi-ronments. Mollisols of tallgrass prairie, like the In-dianterai (Quadeetal.1995),haveasmuchas10wt%organic carbon to depths of more than 1 m, whereascomparable amounts of soil carbon under woodlandsseldomare found deeper than 10 cm (Franzmeier et al.1985). Pedogenic carbonate isotopes of carbon revealentry of C4 grasslands—and deep-calcic Mollisol pa-leosols demonstrate that they were tall grasslands—invading the Himalayan foreland by 7 Ma (fig. 10D).Bunch grasslands had expanded their global climaticand geographic range for the past 34 My (Retallack2013). Accelerated weathering in rangelands in-creased export of nutrient cations to raise phyto-plankton productivity and increase organic matterburial in the ocean, such as the Indus and Bengalsubmarinefans (France-LanordandDerry1997;Guptaet al. 2004). Global cooling was promoted by grass-land expansion because of high grassland albedo(17%–19%),especiallywhencoveredbysnow(albedo1 50%), comparedwith that ofwoodlands (9%–14%)that were replaced by grasslands (Myhre et al. 2005).In addition, grasslands have moist soil but dry air,but woodlands transpire as much water as they canand so have soil drier by 20%–30% than grasslands(Breshears and Barnes 1999). Transpired water vaporwould be a greater problem for global warming thanCO2, were it not so easily rained out. Coevolution ofgrasslands was thus a force for both cooling and dry-ing over geological time, like the evolution of treesushering in the Permian-Carboniferous ice age (Ber-ner 1997) and of early land plants inducing the Hir-nantian ice age (Lenton et al. 2012).

Conclusions

Available evidence of paleosols in India limits theonset of near-modern, highly seasonal, monsoonalprecipitation to between 48 and 20 Ma (fig. 3B). Ourincomplete paleosol time series does not contradictEocene (40 Ma) onset of monsoonal seasonality in-

18 G . J . R E T A L L A CK E T A L .

ferred from fossil wood, snails, and mammals inMyanmar (Licht et al. 2014b, 2015) or climate mod-els suggestingmonsoons as early as 45Ma (Ramsteinet al. 1997; Huber and Goldner 2012). Nonmon-soonal early Eocene climates inferred from paleosolsdocumented here and by Singh et al. (2009) also areapparent from early Eocene fossil floras in southernChina (Spicer et al. 2016). These are older than ra-diometric and structural evidence for uplift of theHimalaya main range beginning at about 23 Ma(Thiede et al. 2004; Clift et al. 2008) and thus lendsupport to the idea that an important threshold indevelopment of near-modern monsoon seasonalityin Indian lowlands was created by the 45–20 Maexpulsion of Paratethyan oceans (Molnar et al. 2010;Bosboom et al. 2014; Carrapa et al. 2015; Patnaik2016) and the 35 Ma uplift to 4 km of the TibetanPlateau (Harris 2006; Rowley and Currie 2006), co-eval with the Oligocene spread of deserts in centralAsia (Sun andWindley 2015; Zheng et al. 2015). Ourevidence of calcic paleosols and low chemical indexof alteration also contradicts the view that Hima-layan uplift created global cooling, culminating inthe Pleistocene ice age (Raymo andRuddiman 1992).Antarctic thermal isolation is an inadequate expla-nation of global cooling for similar reasons of declin-ing Antarctic chemical weathering (Retallack et al.2001). Cenozoic global cooling was more likely dueto the spread of carbon-sequestering grasslands (Re-tallack 2013). Grasslands had been expanding theirrange since the earlyOligocene (34Ma), but tallgrasslands first reached what is now northern India and

NepalbythelateMiocenetoPliocene(7–5Ma;Quadeet al. 1995).Computer modelers of the geological history of

the South Asianmonsoon (Huber andGoldner 2012;Roe et al. 2016) have long lamented the lack of proxydata, particularly on paleoclimatic seasonality, inSouth Asia. Paleosol carbonate metrics now supplynew avenues of proxy data to the question. There isno prospect of gaining comparable data from non-calcareous Cenozoic paleosols of southeast China,Myanmar, andMalaysia (Nichols andUttamo 2005;Biasatti et al. 2012; Licht et al. 2014a), but there isvast potential to expand these results to calcareouspaleosols of India, Tibet, Pakistan, northwest China,and interior Asia (Retallack 1991, 1995; Quade et al.2003; Rowley and Currie 2006; DeCelles et al. 2007;Sun and Windley 2015; Macauley et al. 2016). Ouradmittedly incomplete first application of paleosolBk metrics to the deep time records of the Indianmonsoon is thus just a beginning.

ACKNOW L EDGMENT S

Discussions with A. Sahni, J. Roering, and N. Shel-don greatly improved themanuscript, as did reviewsof previous versions by P. DeCelles and J. Quade. Forassistance during fieldwork in India we thank M.Sharma and S. R. Mishra, and for help in China M.Yu, H. Ge, and X. Gou. Chinese fieldwork wasfunded by National Science Foundation of Chinagrant 41210002 to X. Liu.

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